ORIGINAL PAPER

Small volume andesite magmas and melt–mush interactionsat Ruapehu, New Zealand: evidence from melt inclusions

Geoff Kilgour • Jon Blundy • Kathy Cashman •

Heidy M. Mader

Received: 1 November 2012 / Accepted: 12 April 2013

� Springer-Verlag Berlin Heidelberg 2013

Abstract Historical eruptions from Mt. Ruapehu (New

Zealand) have been small (\0.001 km3 of juvenile magma)

and have often occurred without significant warning.

Developing better modelling tools requires an improved

understanding of the magma storage and transport system

beneath the volcano. Towards that end, we have analysed

the volatile content and major element chemistry of

groundmass glass and phenocryst-hosted melt inclusions in

erupted samples from 1945 to 1996. We find that during this

time period, magma has been stored at depths of *2–9 km,

consistent with inferences from geophysical data. Our data

also show that Ruapehu magmas are relatively H2O-poor

(\2 wt%) and CO2-rich (B1,000 ppm) compared to typical

arc andesites. Surprisingly, we find that melt inclusions are

often more evolved than their transporting melt (as inferred

from groundmass glass compositions). Furthermore, even

eruptions that are separated by less than 2 years exhibit

distinct major element chemistry, which suggests that each

eruption involved magma with a unique ascent history.

From these data, we infer that individual melt batches rise

through, and interact with, crystal mush zones formed by

antecedent magmas. From this perspective, we envision the

magmatic system at Ruapehu as frequently recharged by

small magma inputs that, in turn, cool and crystallise to

varying degrees. Melts that are able to erupt through this

network of crystal mush entrain (to a greater or lesser

extent) exotic crystals. In the extreme case (such as the 1996

eruption), the resulting scoria contain melt inclusion-bear-

ing crystals that are exotic to the transporting magma.

Finally, we suggest that complex interactions between

recharge and antecedent magmas are probably common, but

that the small volumes and short time scales of recharge at

Ruapehu provide a unique window into these processes.

Keywords Andesite � Volatile � Melt inclusions �Ruapehu � Crystal mush � Antecryst � H2O � CO2 �Magma mixing

Introduction

Magma erupted from andesitic volcanoes often records a

complex history, and interactions between recharge and

antecedent magmas and/or crystal mush zones are common

(e.g., Nakamura 1995; Murphy et al. 1998; Nakagawa et al.

1999, 2002; Devine et al. 1998). Most studies that have

highlighted this complexity, however, have focussed on

moderate to large volume andesitic eruptions, where subtle,

fine-scale interactions may be obscured by large recharge

volumes. Mt. Ruapehu, New Zealand, a frequently active

andesite volcano that has historically erupted very small

magma volumes thus provides an interesting test case for

imaging the complexity of subvolcanic magmatic pro-

cesses. In addition, eruptions from Ruapehu have been

Communicated by G. Moore.

Electronic supplementary material The online version of thisarticle (doi:10.1007/s00410-013-0880-7) contains supplementarymaterial, which is available to authorized users.

G. Kilgour � J. Blundy � K. Cashman � H. M. Mader

School of Earth Sciences, University of Bristol, Wills Memorial

Building, Bristol BS8 1RJ, UK

G. Kilgour (&)

Wairakei Research Centre, GNS Science, Taupo 3330,

New Zealand

e-mail: g.kilgour@gns.cri.nz

K. Cashman

Department of Geological Sciences, 1272 University of Oregon,

Eugene, OR 97403, USA

123

Contrib Mineral Petrol

DOI 10.1007/s00410-013-0880-7

extremely difficult to predict (Sherburn et al. 1999), due to

both the small volumes of magma involved in individual

eruptions and the presence of an active hydrothermal sys-

tem, which generates a rather noisy background seismic

signal (Hurst 1998). Therefore, a secondary goal of our

study is to improve our knowledge of both the magma

plumbing and magma transport systems beneath Ruapehu.

Pre-eruptive conditions of magma storage and recharge

can be determined by analysing the compositions of co-

existing crystals (phenocrysts and microlites) and glasses

(groundmass and crystal-hosted melt inclusions). Together

these data provide information on storage temperature and

pressure, magma–magma interactions, and the mixing of

magma and crystal mush zones (e.g., Roedder 1984;

Blundy and Cashman 2005; Liu et al. 2006). When com-

bined with airborne gas chemistry monitoring (e.g.,

Christenson et al. 2010), the volatile content of the glass

phases also provides information on the fate of the main

volatile components during degassing.

We present volatile and major element chemistry of

Ruapehu melt inclusions and groundmass glass from more

than 50 years of eruptions (1945–1996). We use samples of

scoria and lava from every historical magmatic eruption

during this period to track changes in magma composition

through time. Importantly, we provide new evidence that

recent eruptions have been driven by small batches of

recharge magma that have mingled with, entrained crystals

from, and remobilised regions of shallow-stored, anteced-

ent magma.

We analysed representative samples (from GNS Science,

New Zealand rock archives) from six magmatic eruptions at

Ruapehu: 1945, 1969, 1971, 1977, 1995, and 1996. In this

paper, we identify eruptions by year, not by the specific

eruption date. This is particularly relevant for the

1995–1996 eruptive episode, as we were not able to dis-

tinguish between eruptions in September and October 1995

and the June and July 1996 eruptions (c.f. Nakagawa et al.

1999, 2002). Similarly, we were unable to identify a unique

eruption date for the 1945 and 1971 samples. Thus, our

samples are from eruptions (1) occurring between March

and December 1945 [1945]; (2) on 22 June 1969 [1969]; (3)

from April to July 1971 [1971]; (4) on 2 November 1977

[1977]; (5) from September to October 1995 [1995]; and (6)

from June to July 1996 [1996] (Table 1).

Geological background

At 2,797 mASL, Mt. Ruapehu is the largest and most active

andesitic stratovolcano in New Zealand. It is located at the

southern end of the Taupo Volcanic Zone (TVZ) and the

summit area is covered by small permanent glaciers and

snowfields (Fig. 1). Ruapehu has been active for at least

200 ka (Gamble et al. 2003). Holocene activity has

involved a series of overlapping craters (Hackett and

Houghton 1989), with the most recent activity confined to

the southernmost crater, currently occupied by the warm

(15–40 �C) and acidic (0–1 pH) Crater Lake (Hackett,

1985).

Historical activity has consisted primarily of frequent

small phreatic (e.g., Kilgour et al. 2007, 2010) and phre-

atomagmatic eruptions through Crater Lake (Healy et al.

1978; Nairn et al. 1979; Houghton et al. 1987). Larger

magmatic events occurred in 1945 (Oliver 1945; Reed

1945; Beck 1950; Gregg 1960), 1969 (Healy et al. 1978),

1975 (Houghton et al. 1987), and 1995–96 (Houghton et al.

1996; Bryan and Sherburn 1999; Nakagawa et al. 1999,

2002; Johnston et al. 2000). Eruptive activity typically

involves surtseyan jets of lake water and steam accompa-

nied by base surges and ballistic fall-out up to 2 km from

the vent. These events are usually confined to the summit

area (Houghton et al. 1987; Kilgour et al. 2010), but rare

strombolian activity and more widespread sub-plinian to

plinian ash falls also occur (Donoghue 1991; Pardo et al.

2011).

Historical eruption narrative

Volcanic activity at Ruapehu has been observed and

recorded since c. 1850 (Gregg 1960; Hackett and Houghton

1989; Reed 1945). More than 40 eruptions have been

reported since 1945, covering a range of eruptive styles and

sizes (B. Scott pers. comm.). Our samples derive from

magmatic eruptions that ejected juvenile scoria and ash

onto the summit plateau and beyond.

Eruptive activity at Ruapehu between March and July

1945 (Reed 1945) occurred prior to the installation of

volcanic monitoring systems. The 1945 eruption initiated

with explosive magmatic activity that produced high steam

plumes and dispersed ash to c. 200 km from the vent

(Johnston et al. 2000). A series of lava domes were then

constructed and partially destroyed, presumably through

mass wasting and sector collapse (Reed 1945; Beck 1950).

The total volume of erupted magma is estimated at

*0.1 km3 (Johnston et al. 2000).

Between 1945 and the next magmatic eruption in 1969,

a limited seismic network was installed and Crater Lake

temperature measurements and chemical sampling were

initiated. A moderately large eruption (0.9 9 106 m3) on

22 June 1969 ejected older lava lithics, lake sediments, and

*5 vol % juvenile scoria (Healy et al. 1978). We have

calculated the bulk volume of juvenile magma to be

4.5 9 104 m3, with a dense rock equivalent (DRE) of

1.7 9 104 m3. This event was preceded by only limited

seismic precursors and a regular Crater Lake heating–

cooling cycle considered to reflect normal activity (Healy

et al. 1978).

Contrib Mineral Petrol

123

During the 3 months prior to the 1971 eruption, Crater

Lake temperatures rose from ca. 25 to 55 �C, and the

abundance and amplitude of volcanic earthquakes (signal-

ling magma movement) increased (Sherburn et al. 1999).

There also appears to have been an increase in volcanic

earthquakes over a few weeks before first eruption on 3 April

1971. The 1971 eruption was a small, phreatomagmatic

event that was confined to the summit plateau. A phreato-

magmatic eruption on 2 November 1977 was small com-

pared to eruptions of 1945, 1969, and 1975 (samples of the

1975 scoria were not available for this study). The 1977

eruption occurred one to 2 weeks after the temperature of

Crater Lake increased from c. 19 to 30 �C, yet without any

coincident volcanic earthquakes (Sherburn et al. 1999). The

absence of volcanic earthquakes could reflect the small

amount of fresh magma injected to shallow levels, as well as

the rather sparse seismometer network installed at the time.

Although limited field data exist for either eruption, they

appear to have been similar in size to a well-characterised

eruption in 2007 (based on photographs of the summit area

after each eruption). Juvenile scoria was erupted during

those events, and we have again assigned a value of 5 wt%

of the bulk deposit as juvenile material. Therefore, if we take

a similar bulk volume to that of the 2007 event (Kilgour et al.

Table 1 Bulk XRF data of scoria from selected historical eruptions from Ruapehu

Eruption year 1945 1969 1971 1977 1977 1995 1995 1995 1995 1996a

Sample number 1945A 22/5/69-1 1971A 1977A 1977-8 31195B 31195A 71195-04 161195-34 130896

Major elements (wt %)

SiO2 60.16 61.10 58.50 60.44 58.88 57.68 57.90 58.02 61.54 57.57

Al2O3 16.87 15.79 16.24 16.43 16.09 16.30 16.33 16.24 16.58 16.40

Fe2O3 5.80 6.22 6.94 7.65 7.17 7.36 7.33 7.25 6.36 7.43

MnO 0.09 0.09 0.10 0.07 0.10 0.11 0.11 0.11 0.08 0.12

MgO 3.55 4.00 5.04 3.62 4.73 5.45 5.42 5.31 3.61 5.37

CaO 6.02 6.10 7.31 7.08 7.30 7.75 7.74 7.67 6.30 7.87

Na2O 3.47 3.36 3.30 2.55 3.11 3.25 3.25 3.22 2.88 3.30

K2O 1.67 2.15 1.54 1.39 1.48 1.36 1.38 1.43 1.64 1.32

TiO2 0.66 0.65 0.62 0.66 0.64 0.62 0.62 0.62 0.67 0.64

P2O5 0.14 0.16 0.14 0.14 0.14 0.13 0.14 0.14 0.14 0.10

LOI 1.49 0.20 0.05 -0.12 0.20 -0.04 -0.07 -0.08 0.00 0.00

Total 99.92 99.81 99.77 99.92 99.83 99.98 100.15 99.93 99.80 100.00

Trace elements (ppm)

As 4 6 3 2 2 1 2 3 3 4

Ba 384 422 358 353 346 307 320 330 390 352

Ce 27 30 25 39 30 26 21 28 23 22

Cr 48 78 117 97 115 105 113 95 92 114

Cu 64 59 68 86 62 66 68 66 72 72

Ga 20 17 17 18 17 17 16 16 17 17

La 18 14 19 22 \5 \5 11 \5 \5 10

Nb \1 7 \1 \1 \1 \1 \1 \1 \1 4

Ni 22 41 57 50 57 65 62 57 45 59

Pb 11 13 12 12 13 13 12 10 14 10

Rb 60 83 54 51 53 48 48 50 56 47

Sc 19 19 23 22 24 24 28 21 21 26

Sr 280 243 261 252 258 262 263 259 269 264

Th 7 9 5 5 5 5 4 4 4 4

U 2 3 3 2 2 2 \1 2 2 1

V 152 153 176 184 178 182 189 175 173 196

Y 20 21 20 19 18 18 17 19 17 19

Zn 65 61 65 64 66 67 68 68 67 68

Zr 134 154 113 121 115 104 103 105 126 106

a The 1996 sample is taken from Gamble et al. (1999)

Contrib Mineral Petrol

123

2010), the amount of magma erupted in each eruption in

1971 and 1977 is of the order 1 9 104 m3 (DRE).

A period of relative quiescence occurred from 1988 to

1994. However, volcanic tremor (c. 2 Hz) at Ruapehu

rose to, and was maintained at, high levels starting in the

early 1990s. During this time period, Crater Lake heating

cycles were often punctuated by steam explosions,

although there were no signs of significant (magmatic)

eruptive activity (Sherburn et al. 1999). A period of rapid

heating of Crater Lake occurred in November 1994.

During this time, the lake temperature increased from c.

19 to 50 �C in 1 month, and numerous phreatic eruptions

occurred within Crater Lake. However, this activity was

not interpreted as indicative of magmatic injection to

shallow levels (Christenson 2000). The lake heating cycle

appeared to be over by April 1995, but was immediately

followed by another heating event. More phreatic erup-

tions occurred from April to early July. By this time, a

steady increase in the Mg/Cl ratio of the lake water

suggested that magma was being injected to shallow

levels and interacting with the hydrothermal system

(Christenson 2000). Moderate levels of seismicity and

Crater Lake temperatures, combined with and the absence

of deformation determined from theodolite levelling sur-

veys (B. Scott 2011 pers. comm.), indicated that the

amount of magma driving the Mg/Cl ratio changes was

very small. On 18 September 1995, a small phreato-

magmatic eruption occurred with few seismic precursors

(Bryan and Sherburn 1999); following this event, a large

phreatomagmatic eruption developed. A lull in activity

from November 1995 to June 1996 allowed direct mea-

surements of the main fumaroles within the inner crater

(Christenson 2000). Tremor increased to pre-September

1995 levels on 15–16 June 1996 and on 17 June, pul-

sating, phreatomagmatic eruptions graded into a more

continuous eruption, with plumes reaching 8.5 km asl

(Prata and Grant 2001). The 1995–1996 eruptions were

approximately two orders of magnitude larger than the

eruptions of the 1960s and 1970s. Tephra dispersal

mapping suggests DRE volumes for the 11–14 October

1995 and the 17–18 June 1996 eruptions of 3 9 107 m3

and 6 9 106 m3, respectively (Cronin et al. 1998; Fig. 2).

Between 1996 and 2007, Ruapehu remained relatively

quiet except for one small phreatic event on 4 October

2006 (Kilgour et al. 2007; Mordret et al. 2010). Then on 25

September 2007, after *9 min of precursory seismic

a b

Fig. 1 a Location map of Mt Ruapehu at the southern end of the

Taupo Volcanic Zone. b Photograph of Ruapehu’s summit plateau

and Crater Lake. This photograph is taken from the South towards the

North with Ngauruhoe in the background (GNS Science archive

image). Approximate locations for the two main vents (North and

Central) beneath the lake (Christenson et al. 2010) are marked

Contrib Mineral Petrol

123

signals (Jolly et al. 2010), a short-lived phreatic eruption

from Crater Lake created a northerly directed blast that

deposited ballistic blocks and surtseyan jets (Kilgour et al.

2010). Mapping of the deposits within the summit plateau

yielded a total volume of *3 9 105 m3. The presence of

juvenile magma in the 2007 eruption is equivocal and so

we have not included these samples in our study.

The synthesis of data from historical eruptions presented

above shows that recent eruptions of Ruapehu are small but

frequent, and may occur without warning. The volumes of

magma involved in priming the magmatic system for

eruption also appear to be small. Here, we use detailed

compositional analyses of samples from these eruptions, in

the context of this eruptive narrative, to improve our

Fig. 2 Variations in eruption volume, magmatic temperature, and

mineral chemistry between historical eruptions at Ruapehu. Dense

rock equivalent (DRE) volumes for 1945 (Johnston et al. 2000), 1969

(Healy et al. 1978), 1971 and 1977 (assuming a similar volume

erupted in 2007 from Kilgour et al. (2010)), and 1995–1996 (Cronin

et al. 1998). Magmatic temperatures were calculated using the

plagioclase-liquid (open triangles) and clinopyroxene-liquid (filled

circles) geothermometers of Putirka (2008), and Fe–Ti oxides

(crosses) using LePage (2003) for 1969 only. Plagioclase core, rim,

and microlite (Plag C, R, M, respectively) anorthite content (An %),

and clinopyroxene (cpx) and orthopyroxene (opx) magnesian number

(Mg#) are also shown. Plagioclase rim MgO content is expressed as a

range (open rectangles). Average values for plagioclase (An) and

pyroxene (Mg#) are denoted by an ‘‘x’’

Contrib Mineral Petrol

123

understanding of the physical conditions of magma storage,

recharge, and eruption at Ruapehu.

Methods

All samples were analysed for bulk rock chemistry by

X-ray fluorescence (XRF) for major and trace elements at

SpectraChem, Wellington (NZ). Samples were then pre-

pared for the analysis of phase compositions by lightly

crushing individual scoria clasts for each eruption (we

combined two scoria clasts from 1995) and hand-picking

phenocrysts for mounting in epoxy resin onto glass slides.

The preparation of crystal-separate thin sections, rather

than grain mounts, minimised the carbon background from

excess epoxy during secondary ion mass spectrometry

(SIMS) analysis of volatiles. Each crystal separate was

polished to ca. 100 lm thickness to expose melt inclusions.

We obtained back-scattered electron (BSE) images using a

Hitachi S-3500 N SEM at the University of Bristol at

15 kV and at a working distance of *15 mm. These

images were used to map melt inclusions trapped in pla-

gioclase, orthopyroxene, and clinopyroxene phenocrysts.

Approximately, 70 % of the melt inclusions were larger

than 25 lm, the minimum spot size of SIMS analyses.

SIMS analyses of the volatiles dissolved in melt inclu-

sions were carried out on Au-coated grain mounts using a

Cameca IMS-4f ion microprobe at the University of

Edinburgh. We constructed working curves 1H/30Si versus

H2O and 12C/30Si versus CO2 (e.g., Blundy and Cashman

2008) using a total of nine rhyolitic glass standards that

range from 0.15 to 4.1 wt% H2O and 0–2,860 ppm CO2.

Standards were run at three intervals during each day to

account for drift in the analyses. 1H was analysed at low

mass resolution. For 12C, interference from 24Mg2? is

significant at the relatively high MgO contents (2.5 wt%)

of Ruapehu melt inclusions. Separation of the Mg and CO2

spectra thus required us to conduct our SIMS measure-

ments at high mass resolution, and first analyse melt

inclusions for CO2 followed by H2O. We did not analyse

the 1945 lava sample via SIMS due to concerns over H2O

diffusion out of the melt inclusions in such slowly cooled

samples (e.g., Hauri 2002).

Electron probe micro-analysis (EPMA) can damage

hydrous silicate glass. For this reason, we measured the

major element composition of the inclusions by EPMA

after SIMS analysis was completed (e.g., Blundy et al.

2010). EPMA was conducted using a CAMECA SX-100

five-spectrometer wavelength dispersive spectrometry

(WDS) instrument at the University of Bristol. Melt

inclusions and groundmass glasses were analysed using a

15-kV accelerating voltage, 4 nA beam current with a

defocused 10 lm beam, with K and Na being analysed first

to reduce the effects of alkali migration (e.g., Humphreys

et al. 2006). Plagioclase and pyroxene crystals were ana-

lysed using a 20-kV accelerating voltage, 10 nA beam

current and a focussed beam. Calibration used a selection

of mineral and oxide standards. Data were reduced using

the ZAF procedure.

Results

Petrography of Ruapehu samples

Samples available for this study include a lava sample from

1945 and scoria from 1969, 1971, 1977, 1995, and 1996.

The scoriae are vesicular, porphyritic, and microlite-rich

(with the exception of 1969), similar to pre-historic Ru-

apehu scoria (Hackett 1985; Graham and Hackett 1987;

Gamble et al. 1999). Here, we use the term phenocryst to

signify a crystal that is significantly larger (*1–2 mm)

than microlites present in the groundmass, independent of

its origin. We use antecryst to mean a crystal that

demonstrably grew within a different magma than its cur-

rent host (i.e., exotic), and cognate to mean a crystal that

grew from and erupted with its host magma.

Phenocrysts of plagioclase dominate the mineral

assemblage with lesser amounts of clinopyroxene and

orthopyroxene (Table 2). However, there is no clear crys-

tallisation sequence; pyroxene is often found within pla-

gioclase phenocrysts, while plagioclase inclusions are also

seen in pyroxene crystals. The implication is that these

three phases precipitated cotectically. Of the minor phases,

magnetite is present in all samples as small blocky grains

up to 30 lm across. Ilmenite is absent in all samples except

in 1969 scoria, in agreement with previous observations

(Nakagawa et al. 1999; Price et al. 2012). Hornblende is

absent in all samples; this is also the case for all but one

lava exposed on the edifice of the volcano (Hackett 1985).

Microlites (crystals \100 lm across) of plagioclase,

clinopyroxene, orthopyroxene, magnetite, and rare ilmenite

Table 2 Representative crystallinity of Ruapehu scoria

Eruption date 1945 1969 1971 1977 1995 1996a

Sample number 1945F 1969A 1971A 1977A 031195B 57536

Phenocrysts (%) 40 29 33 31 26 36

Groundmass (%) 59 42 50 49 54 39

Vesicle (%) 1 29 17 20 21 25

Phenocrysts

Plagioclase 29 20 23 22 18 19

Clinopyroxene 6 4 5 4 3 8

Orthopyroxene 5 5 5 4 5 9

a The 1996 sample is taken from Gamble et al. (1999)

Contrib Mineral Petrol

123

are present in the groundmass of most samples (c.f. Nak-

agawa et al. 1999).

The bulk compositions of scoria and lava samples from

1945 to 1996 have been reported by Gamble et al. (1999)

and Nakagawa et al. (1999) and pre-historical compositions

by Price et al. (2012). We have conducted further bulk rock

XRF analyses on samples from historical eruptions, which

agree well with data from Gamble et al. (1999) (Table 1).

There are no discernible systematic changes in bulk

chemistry, which ranges from *57 to 64 wt% SiO2,

mineral content or abundance, or isotopic composition

through time (Gamble et al. 1999). This chemical monot-

ony attests to a genetic link among all Ruapehu magmas.

Interestingly, scoria erupted during the 1995–1996 erup-

tions span the entire range of the historical record

(*58–62 wt% SiO2).

Mineralogy

Plagioclase phenocrysts are up to 4 mm across and are

generally zoned from a calcic core (An55–82) to a sodic rim

(An52–65) (Fig. 2). The anorthite (An) content of pheno-

cryst rims is similar to that of plagioclase microlites

(An50–65). Rare plagioclase phenocrysts are reversely

zoned from *An52 to * An60. The MgO content of pla-

gioclase is commonly used to monitor mafic inputs to the

magmatic system (e.g. Hattori and Sato 1996). There is not

much variation in MgO content among our samples, except

for 1971 and 1995 samples, which show elevated MgO in

some plagioclase phenocryst rims (up to 0.15 wt%; see

also Nakagawa et al. 1999).

Clinopyroxene phenocrysts are subhedral to euhedral

and B3 mm in length. All analysed samples show the same

compositional range (with a mean of Wo*42; En*43;

Fs*15; Fig. 2). Oscillatory zoning of clinopyroxenes is

evident in all Ruapehu scoria (e.g., Nakagawa et al. 1999,

2002). Normal zoning (Mg#53–77) is most common,

although many clinopyroxene crystals (both phenocrysts

and microlites) preserve a very thin (2–5 lm) Mg-rich

outermost rim (Fig. 3). Melt inclusions are common as

both small (\10 lm), glassy inclusions forming concen-

trically to the growth pattern and large ([50 lm), isolated

inclusions near the core of the crystal.

Orthopyroxene phenocrysts are euhedral to subhedral,

relatively unzoned, and B4 mm in length, with a compo-

sition of Wo*3; En*60; Fs*37—enstatite (Fig. 2). Rare

zoned (both normal and reverse; Mg#36–50) orthopyroxene

crystals are found in all historical samples. Orthopyroxene-

hosted melt inclusions are less common than in clinopy-

roxene and are present as small (\ 10 lm) inclusions on

the margins of the crystal. Larger ([30 lm) melt inclusions

are uncommon in phenocrysts from all eruptions.

Major element chemistry of groundmass glass

The groundmass glasses of Ruapehu scoria span a wide

compositional range from 58 to 78 wt% SiO2 (Table 3;

Figs. 4a, 5, Supplementary Table 1). Glass compositions

from individual Ruapehu eruptions can be distinguished

from each other by means of major element binary plots.

The least evolved glasses are found in 1995 and 1996

scoria (Fig. 5). Groundmass glasses from 1945 lava and

1969 scoria are significantly more evolved than all other

historical samples; in the case of 1945, this is possibly due

to slower cooling of the lava sample. Conversely, glasses

from the 1971 and 1977 eruptions extend to lower SiO2

a b

Fig. 3 a BSE image of a clinopyroxene phenocryst from the 1995

eruption of Ruapehu. Oscillatory zoning is common throughout all

eruptions. The greyscale images highlight zones of relatively higher

Fe (light grey) and Mg (dark grey) zones. Note the more diffuse

boundary at the core compared to the sharp boundary at the rim.

b Dashed black lines denote the boundary between relatively diffuse

Fe- and Mg-rich zones. Arrow points to the *3 lm wide, dark greyrim (Mg-rich) on the outermost margin of the crystal, elsewhere

interpreted as late-stage mixing (e.g., Saunders et al. 2012). Ground-

mass glass (gl) is shown in (b)

Contrib Mineral Petrol

123

with little overlap in major element composition with the

1969 glasses. This offset is most clearly seen in Na2O

(Fig. 4a). Therefore, although some of the eruptions ana-

lysed here were less than 2 years apart (1969 and 1971)

their groundmass glass composition is distinct (Fig. 4a–d).

A further striking feature of the groundmass glasses is

the wide range in SiO2 content (4–8 wt%) within an indi-

vidual eruption, attesting to the heterogeneous nature of

groundmass glasses from Ruapehu (Fig. 4a–d). For

instance, the SiO2 content of groundmass glasses erupted in

1969 ranges from 69 to 74 wt%, 1971 from 65 to 73 wt%,

1977 from 66 to 70 wt%, and 1995–96 from 59 to 66 wt%.

The latter are less evolved, however, than the range of

62–70 wt% SiO2 reported for the 1995–1996 eruptions by

Nakagawa et al. (1999); we are not able to explain this

discrepancy. For all eruptions, the groundmass glass

exhibits a linear trend towards the bulk rock composition in

most binary major element plots (Fig. 4a–d). A notable

exception is Na2O (Fig. 4a), where linear trends of nega-

tive slope do not extrapolate to the bulk rock composition.

Major element chemistry of melt inclusions

The major element chemistry of some pyroxene- and

plagioclase-hosted melt inclusions is presented in Table 4

(for the full dataset, refer to Supplementary Table 2) and

plotted in Fig. 4e–h. The major element composition of

pyroxene- and plagioclase-hosted inclusions is similar in

all samples and covers the same range as the groundmass

glass (i.e., 60 to 73 wt% SiO2). Unlike the groundmass

glass compositions, however, plagioclase- and pyroxene-

hosted melt inclusions from individual eruptions do not

form distinct clusters in major element plots, but instead

exhibit significant overlap between eruptions (Fig. 4e–h).

Overall, however, the most evolved inclusions are from

the earliest eruption analysed (1945), and the least

evolved inclusions are from the 1995–1996 eruptions

(Fig. 5).

Volatile content of melt inclusions

Our SIMS analyses (CO2, H2O, Li, Be, B, F, Cl) of melt

inclusions from historical eruptions at Ruapehu (Table 4;

Fig. 6) are the first direct measurements of the volatile

content of Ruapehu magmas. Most Ruapehu melt inclu-

sions have H2O contents of 1–1.5 wt%, with a small

number of inclusions exceeding 2.5 wt% (Fig. 7), which is

relatively dry compared to similar andesitic systems in arc

settings (e.g., Blundy et al. 2010; Devine et al. 1998;

Portnyagin et al. 2007; Wallace 2005). According to the

literature, the only arc andesite magma with a similarly low

H2O content is from the 1994 to 1998 eruption of Popo-

catepetl (Mexico), with a range of 0.8–3.02 wt% H2O

(Atlas et al. 2006). The low H2O contents have implica-

tions for magma evolution, phase relations, and transport

properties.

Ruapehu melt inclusions range in CO2 from c. 25 to

1,059 ppm, with most inclusions having less than c.

600 ppm CO2 (Table 4; Fig. 7). These CO2 values are

significantly higher than most intermediate subduction

zone magmas (Wallace 2005) but are again similar to those

from Popocatepetl, where melt inclusions preserve

B1,458 ppm CO2 (Atlas et al. 2006).

Melt inclusions preserve elevated halogen contents with

up to 2,069 ppm F and 1,342 ppm Cl (Table 4; Fig. 6d, g,

h). Average halogen values in melt inclusions are c.

955 ± 175 ppm F and 659 ± 128 ppm Cl; average

groundmass glass values are 838 ± 138 ppm F and

515 ± 88 ppm Cl. These data suggest only limited

degassing at very low pressures, consistent with experi-

ments on basaltic bulk compositions that show that the

significant Cl loss requires very low pressures, after most

Table 3 Average major element composition of groundmass glass from historical Ruapehu eruptions

Major Element (wt %) 1945 1969 1971 1977 1995 1996

SiO2 77.28 (0.91) 72.07 (1.24) 68.88 (2.08) 67.93 (1.07) 63.2 (1.24) 62.06 (1.23)

Al2O3 12.41 (0.7) 13.32 (0.91) 14.36 (1.34) 13.95 (0.43) 15.34 (0.59) 15.68 (1.2)

FeO 1.82 (0.35) 3.2 (0.52) 5.09 (0.89) 6.07 (0.39) 6.47 (0.55) 4.61 (3.24)

MnO 0.02 (0.01) 0.05 (0.02) 0.07 (0.02) 0.09 (0.02) 0.11 (0.04) 2.48 (3.37)

MgO 0.11 (0.13) 0.49 (0.32) 0.82 (0.53) 1.22 (0.17) 2.5 (0.61) 2.45 (0.48)

CaO 1.03 (0.37) 2.06 (0.51) 3.45 (0.69) 3.9 (0.3) 5.58 (0.6) 5.85 (0.66)

Na2O 1.97 (0.42) 3.83 (0.31) 2.88 (0.84) 2.33 (0.67) 3.18 (0.61) 3.36 (0.65)

K2O 4.47 (0.45) 4.15 (0.3) 3.01 (0.31) 3.05 (0.16) 2.45 (0.26) 2.36 (0.45)

TiO2 0.88 (0.23) 0.68 (0.08) 1.17 (0.18) 1.19 (0.05) 0.99 (0.08) 1.01 (0.11)

P2O5 0.19 (0.05) 0.13 (0.05) 0.25 (0.06) 0.25 (0.04) 0.17 (0.08) 0.2 (0.04)

Number of analyses 18 36 17 31 28 45

Standard deviations are given in parentheses. The full dataset can be found in Supplementary Table 1

Contrib Mineral Petrol

123

of the initial H2O, SO2, and CO2 has already exsolved

(Lesne et al. 2012).

Li concentrations in melt inclusions vary between 22 and

80 ppm, with most of the data confined to between 40 and

70 ppm for all eruptions (Table 4; Fig. 6). Matrix glasses

contain 26–50 ppm Li. Li contents of melt inclusions and

groundmass glasses tend to increase with increasing H2O

(Fig. 6a), suggesting that Li partitions modestly into the

vapour phase during degassing. Be concentrations are low,

from 1 to 5 ppm, and show no correlation with H2O

(Fig. 6b). Be concentrations in the matrix glass are similar to

the lowest value in melt inclusions of *1 ppm. B contents

a e

b f

c g

d h

Fig. 4 Major element plots of groundmass glass (a–d) and pheno-

cryst-hosted melt inclusion compositions (e–h) from 1945 to 1996.

Bulk rock XRF data are shown as the open black ellipses. All data are

re-calculated to anhydrous values. Note the distinction between each

eruption is clear in the groundmass glass, yet the melt inclusions are

relatively tightly clustered with significant overlap. Samples are from

the same suite of samples as Fig. 2 and Table 1

Contrib Mineral Petrol

123

of most melt inclusions range from *30 to *80 ppm and

are not correlated with H2O (Fig. 6c). Matrix glasses range

from *30 to *50 ppm B, with some higher values

recorded in groundmass glass from the 1969 eruption. Evi-

dently, the behaviour of B and H2O is decoupled during

magmatic degassing.

Fig. 5 Plot of K2O versus SiO2

showing the evolution of

groundmass glasses and melt

inclusions to less-evolved

compositions with time. All

data have been re-calculated to

anhydrous values. Symbols are

the same as in Fig. 4

Contrib Mineral Petrol

123

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Contrib Mineral Petrol

123

Ta

ble

4co

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Mel

tin

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Sec

on

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ple

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(wt%

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1.1

90

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0.0

91

00

71

50

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Py

rox

ene

19

77

py

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inc

02

1.4

26

24

34

63

97

08

49

66

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1.0

43

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12

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5.7

73

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42

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0.0

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08

44

0.3

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py

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03

1.4

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42

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01

01

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0.9

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rox

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03

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001

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rox

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py

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rox

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rox

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Contrib Mineral Petrol

123

Ta

ble

4co

nti

nu

ed

Mel

tin

clu

sio

ns

Sec

on

dar

yio

nm

ass

spec

tro

met

ry(S

IMS

)d

ata

Ele

ctro

np

rob

em

icro

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aly

sis

(EP

MA

)d

ata

(wt%

)P

ress

ure

(bar

s)

XH

2O

Ho

st

min

eral

Sam

ple

nu

mb

erH

2O

(wt%

)

CO

2

pp

m

Li

pp

m

Be

pp

m

B pp

m

F pp

m

Cl

pp

m

SiO

2T

iO2

Na 2

OA

l 2O

3F

eOK

2O

CaO

Mg

OC

r 2O

3M

nO

To

tal

19

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py

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inc

02

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0

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21

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19

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23

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36

0.7

91

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07

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2.9

05

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2.8

50

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0.1

61

00

Contrib Mineral Petrol

123

Discussion

Magmatic storage conditions

We have determined magmatic temperatures for the his-

torical Ruapehu eruptions using several different geother-

mometers. The absence of touching Fe–Ti oxide pairs in all

samples (except for disparate pairs in the 1969 sample)

precludes us from using the method of Lindsley and

Anderson (1983); for this reason, we have used the pla-

gioclase-liquid (Putirka 2008), clinopyroxene-liquid (Put-

irka 2008), orthopyroxene-liquid (Putirka 2008), and two-

pyroxene (Lindsley and Frost 1992) geothermometers

(assuming a H2O content of *1.5 wt%). In all geother-

mometer calculations, we have assumed a pressure of

250 MPa, which is based on the volatile data (see below)

and geophysical interpretations (Ingham et al. 2009; Row-

lands et al. 2005). While these methods yield different

absolute temperatures, the trend in the data from one

eruption to another is consistent (Fig. 2). All historical

eruptions plot between 910 and 1,080 �C, with the bulk of

the data clustering between 950 and 1,050 �C. There is a

general increase in magmatic temperature from the earliest

sample analysed (1945) to the latest scoria sample (1996)

(Fig. 2). The relatively low temperature for 1945 may

reflect slow cooling and re-equilibration of this lava sample.

Temperature estimates for the 1995–1996 eruptions

occupy a limited range between 1,000 and 1,080 �C with

no apparent clustering towards the high or low estimates.

These data contrast with those of Nakagawa et al. (1999),

who found two separate populations of clinopyroxene–

orthopyroxene pairs, one that yielded temperatures of

a e

b f

c g

d h

Fig. 6 Trace element (Li, Be,

B, Cl) variation with H2O

measured by SIMS. a Li

increases with increasing H2O.

Be b and B c show no

correlation with H2O. d A weak,

positive correlation exists

between Cl and H2O. (e–h)

Trace element (Li, B, F, Cl)

content versus pressure plots. Cl

degasses at low pressure, while

the other trace elements show

no change with pressure

Contrib Mineral Petrol

123

*1,000 �C and another with temperatures of 1,000–

1,200 �C. The high temperatures, which were calculated

using the method of Lindsley and Anderson (1983), appear

unreasonably high for the andesite composition of the

Ruapehu ejecta. For instance, the liquidus temperature of

the 1995–1996 magma is approximately 1,150 �C (using

Danyushevsky and Plechov, 2011) which constrains the

maximum phenocryst temperature. Also, the updated geo-

thermometers of Putirka (2008) produce a more homoge-

neous temperature than the Lindsley and Anderson (1983)

method. Therefore, for the purposes of this paper, we have

used the Putirka (2008) geothermometers throughout.

The volatile contents of melt inclusions are similar for

all analysed eruptions (Figs. 6, 7). They contain

B1,000 ppm CO2 and *1.5 wt% H2O, which suggests a

minimum (volatile saturation) trapping pressure of

*50–270 MPa (at between 920 and 1,030 �C using the

calculation of Papale et al. 2006). This is in agreement with

estimates from the phenocryst melt and two-pyroxene

geobarometers of Putirka (2008). If we assume a crustal

density of 2,600 kg/m3 and volatile saturation, this pressure

range suggests that the magma storage region beneath

Ruapehu extends from *2 to 9 km below the volcano.

Our data show that Ruapehu magmas are relatively dry

compared to other arc andesites. While the phase equilibria

of Moore and Carmichael (1998) appear consistent with a

dry, andesitic magma, it has been proposed that pheno-

cryst-hosted melt inclusions are able to rapidly hydrate or

a

b

Fig. 7 Volatile contents as measured by SIMS against modelled

degassing histories for pyroxene-hosted melt inclusions from Ru-

apehu. a H2O versus CO2 content of melt inclusions plotted with

calculated isobars (dashed grey lines) and vapour isopleths in mol %

H2O (dashed black lines). Isopleths and isobars were calculated using

Papale et al. (2006). Illustrative closed-system degassing curves

(curved black lines) are plotted from three starting compositions

(Curve 1–1.28 wt% H2O and 1,060 ppm CO2, Curve 2–1.75 wt%

H2O and 800 ppm CO2, and Curve 3–2.18 wt% H2O and 700 ppm

CO2). b Plot of calculated XH2O versus saturation pressure showing

the coherence of all melt inclusions between closed-system degassing

paths 1 and 3. Significant H2O-loss or CO2-fluxing would result in the

inclusion population following a systematic decrease in XH2O, with a

slight decrease in saturation pressure (arrowed line). Rare inclusions

that fall below degassing curve 1 in (b) probably did lose significant

H2O. Closed-system degassing of distinct magmas with a similar

volatile content can explain the observed variations in H2O versus

CO2 and XH2O versus pressure space. Average propagated errors in

the SIMS analyses are shown

Contrib Mineral Petrol

123

de-hydrate due to H? diffusion (e.g., Gaetani et al. 2012)

and rapid equilibration with the surrounding melt. Lloyd

et al. (2013) also suggest that melt inclusions within large

scoria or lapilli clasts are prone to significant dehydration.

Clearly, the H2O content of our melt inclusions will have a

significant effect on the pressure determinations. In order to

assess our melt inclusion measurements, we plotted H2O

versus CO2 (Fig. 7a) and the calculated XH2O (mol frac-

tion H2O of the vapour) against saturation pressure

(Fig. 7b). We then compared our data to modelled open or

closed-system degassing profiles (Papale et al. 2006). Our

data are weakly scattered around a mean of 1.5 wt% H2O

and range from 50 to 1,000 ppm CO2. While these data

could record a complex interplay between CO2-fluxing,

crystallisation, and H2O loss (e.g., Blundy and Cashman,

2008, Spilliaert et al. 2006), we consider a simpler inter-

pretation. In the XH2Ovapor versus pressure diagram, all the

data appear to follow a relatively simple closed-system

degassing profile wherein XH2O increases with decreasing

pressure (Fig. 7b). There is some scatter in the data, which

may reflect different initial volatile contents or degassing

trajectories. However, the data are not consistent with

significant diffusive loss of H2O from the melt inclusions.

If the inclusions had dehydrated significantly, we would

expect a large number (if not all) of the melt inclusions to

record low XH2Ovapor and anomalously low pressures,

inconsistent with other independent data. To illustrate the

effect of diffusive H2O loss, we took one melt inclusion

composition and progressively reduced its H2O content by

0.5 wt%, from 2.5 wt%. At each point, we calculated the

XH2O and saturation pressure. The result provides a vector

for which melt inclusions would trend towards given sig-

nificant H2O loss (Fig. 7b). From this, we can see that

Ruapehu melt inclusions do not exhibit significant H2O

loss.

To explain the H2O and CO2 data, we first must consider

that the bulk chemistry of Ruapehu magmas is broadly

similar (Table 1), with relatively small variations between

eruptions. This indicates that all eruptions are derived from

a similar parental magmatic system (e.g., Gamble et al.

1999). Therefore, we considered the degassing trajectories

of three magmas with broadly similar major element

chemistry and variable initial H2O and CO2. Most of the

data plot between the two bounding closed-system degas-

sing curves (curve 1–1.28 wt% H2O and 1,060 ppm CO2;

curve 3–2.18 wt% H2O and 700 ppm CO2). Based on these

data, we can conclude that Ruapehu magmas have a H2O

content of up to 2.18 wt% and a CO2 content of at least

700 ppm.

As stated above, ilmenite is only observed within scoria

from 1969 (magnetite is noted in all scoria). In order to

calculate equilibrium magmatic temperature and oxygen

fugacity, it is best to analyse touching ilmenite–magnetite

pairs. However, as we were unable to find touching pairs in

the 1969 sample, we tested all possible combinations of

disparate ilmenite and magnetite compositions for equi-

librium using the method of Bacon and Hirschmann

(1988). Only equilibrium pairs were used to determine the

fugacity and temperature of the 1969 magma using ILMAT

(LePage 2003). We calculated an oxygen fugacity of log

fO2 -11.20, equivalent to the Ni-NiO oxygen buffer (NNO)

and temperature c. 939 �C (O’Neill and Pownceby 1993;

Frost 1991), which is *20 �C higher than the average of

the plagioclase-liquid and pyroxene-liquid geothermome-

ters. The remarkably similar mineralogy and phenocryst

compositions suggest that the oxygen fugacity of all his-

torical Ruapehu magmas lies close to NNO, similar to pre-

historical magmas (Price et al. 2012).

To further assess the consistency of our pressure–tem-

perature estimates, we compare our data to high tempera-

ture and pressure melting experiments on andesites. The

bulk composition of Ruapehu magmas is similar to that of

Volcan Colima, Mexico (Moore and Carmichael 1998),

thus the starting compositions for hydrous phase equilibria

experiments from that volcano provide a reasonable com-

parison to natural Ruapehu samples, except that the Colima

experiments were run under H2O-saturated conditions, that

is PH2O = Ptot. Because Ruapehu magmas are H2O-poor

(PH2O \ Ptot), a pressure correction must be applied to the

experimental data. This correction is relatively straight-

forward because of the negligible effects of CO2 on phase

equilibria in silicate systems at low pressures. If we take

the most volatile-rich melt inclusion of *1,000 ppm CO2

and 1.5 wt% H2O, we calculate a saturation pressure of

*270 MPa, with an XH2O of *12–16 % at a temperature

range of between 915 and 1,030 �C (Papale et al. 2006).

This equates to a PH2O of *32–43 MPa. If we assume that

the addition of CO2 simply increases the Ptot, without

effecting phase relations, then we can use this value of

PH2O to match our data to the experiments. Using this

correction, Ruapehu magmas plot near the 2 wt% H2O

isopleth, in a region mostly outside of the hornblende

stability field, but with plagioclase, orthopyroxene, clino-

pyroxene, and magnetite stable, in accord with the

observed phenocryst populations. The experimental equi-

librium plagioclase composition is *An60–65, consistent

with the measured composition of plagioclase rims

(Fig. 8). From this comparison, we conclude that the

mineralogy, volatile content, and temperature of Ruapehu

magma are consistent with the andesite phase equilibria of

Moore and Carmichael (1998) at PH2O of *40 MPa.

Interaction between magma and crystal mush

A compositional comparison of melt inclusions, ground-

mass glass, and bulk rock can be used to determine the

Contrib Mineral Petrol

123

extent to which melt inclusions and their phenocryst hosts

are in chemical equilibrium with the host magma. How-

ever, we must first consider whether any of the melt

inclusions have been modified by post-entrapment crys-

tallisation. Daughter minerals are absent from all melt

inclusions analysed. To evaluate the extent of melt inclu-

sion crystallisation onto the host crystal, we plotted

separately MgO versus Al2O3 (Fig. 9) for plagioclase- and

pyroxene-hosted inclusions, as these elements are differ-

ently compatible in pyroxene and plagioclase crystals. For

example, pyroxene-hosted melt inclusions that crystallise

on the host would result in a displacement to very low

MgO content with little change in Al2O3, whereas crys-

tallisation of plagioclase would lead to a decrease in Al2O3

and slight increase in MgO. These different trends are

shown as vectors corresponding to 5 wt% crystallisation in

Fig. 9. In general, the compositional overlap between

plagioclase- and pyroxene-hosted melt inclusions suggests

that post-entrapment crystallisation was limited. Specifi-

cally, compositional variations in the melt inclusions fol-

low cotectic crystallisation trends; that is, the trends do not

follow the vectors anticipated for post-entrapment crystal-

lisation of the host mineral (Fig. 9).

As melt inclusions do not appear to have experienced

significant post-entrapment crystallisation, they can be

used to examine variations in crystallisation (driven by

cooling, decompression, and/or H2O loss) and magma

mingling/mixing in the small magma batches produced by

Ruapehu eruptions. It is useful to discuss these data in two

groups: (1) 1945–1977 and (2) 1995–1996. We use the

incompatible elements K2O and TiO2 as plotting parame-

ters as these two components best illustrate the composi-

tional differences between eruptions.

1945 to 1977

The major element chemistry of plagioclase- and pyrox-

ene-hosted melt inclusions from 1945 plots along the same

fractional crystallisation trend as the bulk rocks (Fig. 10).

The melt inclusion compositions occupy a wider range than

the groundmass glass. Under equilibrium conditions, the

compositions of melt inclusions should lie between the

bulk rock and the groundmass glass on a plot of two

Fig. 8 Phase diagram of an andesite of similar bulk composition to

Ruapehu (Volcan Colima), in terms of PH2O versus temperature,

redrawn from Moore and Carmichael (1998). Grey box represents the

calculated (using Papale et al. 2006) PH2O conditions of H2O-poor,

CO2-rich historical Ruapehu magmas. Magmatic temperatures were

determined by crystal-melt and Fe–Ti oxide geothermometry.

Ruapehu magmas occupy a PH2O-temperature space where the

equilibrium phase assemblage consists of plagioclase (Plag), ortho-

pyroxene (Opx), clinopyroxene (Aug), and magnetite (Mt). Horn-

blende (Hbl) is absent from Ruapehu due to the relatively low H2O-

saturated pressure and high-temperature conditions of Ruapehu

magmas. Sub-horizontal dashed lines are H2O concentration isopleths

Fig. 9 MgO v Al2O3 plot of

plagioclase- and pyroxene-

hosted melt inclusions from the

1995 eruption of Ruapehu.

These data show that the

chemistry of each inclusion is

largely independent of the host

mineral. Plagioclase- and

pyroxene-hosted melt inclusions

that had crystallised after being

trapped would trend away from

their respective host crystal

along the vectors shown as

black arrows. The length of

arrows approximates 5 %

crystallisation of plagioclase

(Plag) and pyroxene (Pyrox)

Contrib Mineral Petrol

123

nominally incompatible major elements, in this case K2O

and TiO2 (e.g., Faure and Schiano 2005). Melt inclusions

that fall outside of the equilibrium line may indicate that

the crystals are exotic to the host magma, as appears to be

the case for the 1945 data. The most likely source for

exotic crystals at Ruapehu is crystal mush, such as that

invoked by Nakagawa et al. (1999), because of both

numerous crystal clots within scoria clasts and the presence

of a high- and low-temperature signature from two-

pyroxene geothermometry. Together, these data suggest

that the 1945 magma probably intersected, and interacted

with, at least one crystal mush zone during ascent.

The 1969 melt inclusions and groundmass glasses plot

along a fractional crystallisation trend from the bulk rock

composition (Fig. 10). Most inclusions are less evolved

than the groundmass glass, which is to be expected if

crystallisation continues in the melt after inclusions

become trapped within a crystal. This suggests that the

1969 eruptions were driven by a small volume magma that

was isolated from the larger magma storage region. In

contrast, most melt inclusions from the 1971 eruption

describe a fractionation trend that is different from the

groundmass glass. The shift of exotic inclusions to high

K2O at constant TiO2 requires that they crystallised from a

Fig. 10 Plots of TiO2 versus K2O (two incompatible elements)

showing the composition of groundmass glass, melt inclusion and

bulk rock XRF data for each Ruapehu eruption analysed. Cognate

melt inclusions should fall on the same line (grey dashed line) as the

bulk rock and groundmass glass (e.g., Faure and Schiano, 2005). The

1945 melt inclusions span a similar range in K2O to the groundmass

glass, while the melt inclusions are displaced to lower TiO2. The

crystals are therefore equivocally exotic. The 1969 melt inclusions are

mostly less evolved than the groundmass glass, appear on a similar

chemical trend to the whole rock, and we conclude that most melt

inclusions (and hence the crystals) are cognate with the groundmass

glass. The 1971 and 1977 melt inclusions are more evolved than the

groundmass and are displaced either side of their respective

equilibrium mixing lines from the groundmass glass and whole rock

compositions. Most of the 1971 and 1977 inclusions are considered

exotic to the host melt. The 1995 melt inclusions have a larger

compositional spread than the groundmass glass. Many inclusions are

of a similar composition to the melt (cognate), while the more mafic

and silicic end members are possibly exotic. The 1996 groundmass

glass is less evolved than the melt inclusions. This implies that all of

the 1996 melt inclusions and hence the entire population of inclusion-

bearing phenocrysts are exotic. The groundmass glass and melt

inclusions are chemically distinct, but appear on the same mixing

line; therefore, both melts must have had a similar parent composition

and mineralogy. In all plots, black triangles represent the bulk rock

composition for scoria from 1945 to 1996 (Gamble et al. 1999; this

work)

Contrib Mineral Petrol

123

melt with a bulk composition that was distinct from the

transporting magma (represented by the groundmass glass).

Importantly, both exotic and cognate melt inclusions can

be found within the same crystal, and without an obvious

spatial distribution. However, there is evidence for some

crystallisation within the transporting magma in the small

number of 1971 inclusions that are chemically similar to

the groundmass glass and could thus be genetically related

to the host magma (i.e., cognate). The lack of overlap

between either the melt inclusion populations or the

groundmass glass compositions in 1971 and 1969 magmas

is significant given the limited time between eruptions.

These data suggest that prior to the eruption, the 1971

magma entrained crystals from a mush zone that was dis-

tinct from the 1969 magma. The exact date or size of

eruptions that provided the antecrysts cannot be con-

strained, but we can discount crystals generated in 1945

based on their different melt inclusion compositions

(Figs. 4, 5).

As seen in the 1971 scoria, most of the phenocryst-

hosted melt inclusions in the 1977 scoria are displaced to a

higher K2O content than the host glass and are therefore

considered exotic (Fig. 10). Moreover, exotic melt inclu-

sions from the 1971 and 1977 eruptions are not chemically

similar, which suggests that the 1971 and 1977 magmas

ascended through different parts of the subvolcanic system.

As also seen in 1971, a small number of 1977 melt inclu-

sions have a composition that is similar to that of the bulk

rock. This suggests that less-evolved magma was intro-

duced into the base of the andesitic magma storage region

shortly before each eruption. Moreover, although the

groundmass glass compositions of 1971 and 1977 scoria

are similar, they are not identical (Fig. 4), and thus suggest

that the two events involved two distinct magmas despite

their extremely small volumes (e.g., Houghton et al. 1987).

1995 to 1996

The groundmass glass and melt inclusions of the 1995

magma span a similar range in major element chemistry to

earlier eruptions (Figs. 4, 5). The bulk of the melt inclu-

sions from 1995 are more evolved than the groundmass

glass, but appear to lie on a similar liquid line of descent,

indicating that some of the crystals are probably cognate.

The groundmass glass and melt inclusion compositions

from 1995 plot near the bulk rock composition (Fig. 10),

which suggests minimal crystallisation prior to eruption.

Pyroxene- and plagioclase-hosted inclusions are both

more- and less-evolved than the groundmass glass. Those

that are more evolved are likely to be hosted by antecrysts

derived from a crystal mush that is genetically linked to the

historical eruptions. In contrast, a striking aspect of the

Fig. 11 Conceptual model for the magmatic system at Ruapehu.

Small volume andesitic melts are residing as sills and dykes from\2

to *9 km depth, based on the volatile content of phenocryst-hosted

melt inclusions (using Papale et al. 2006), combined with magneto-

telluric soundings (Ingham et al. 2009) and seismic tomography

(Rowlands et al. 2005). A hydrothermally altered boundary zone

(grey diffuse boundary) exists on the margins of the dyke system

(Ingham et al. 2009). Small volume sills and dykes are distinct but

closely spaced. When magma ascends, it interacts with a crystal mush

zone/s (1 and 2), entraining some of those exotic crystals into the

melt. On ascent, small defects in individual crystals allow for the

ingress and then trapping of cognate melt inclusions (3). These

cognate melt inclusions record the magmatic conditions of the new

melt alongside those that of previous melt/s

Contrib Mineral Petrol

123

1996 scoria is that most of the melt inclusions are more

evolved than the hosting glass, which is similar in com-

position to that of the 1995 eruption (Fig. 10). For this

reason, we consider that most, possibly all, of the melt

inclusions analysed from the 1996 eruption are exotic to

the transporting melt. More-evolved melt inclusion com-

positions may indicate that the mush from which antecrysts

were entrained had cooled/crystallised significantly prior to

interaction.

There is substantial overlap in the compositions of

plagioclase- and pyroxene-hosted inclusions from the 1995

and 1996 scoria; because some of the 1995 crystals are

considered cognate, it seems likely that some of the

phenocrysts from 1996 are antecrysts that originally grew

in the 1995 magma (Fig. 10). That the groundmass glass

from 1996 is distinctly more mafic than the melt inclusions

suggests that the 1996 eruption involved a very crystal-

poor, relatively mafic magma that entrained crystals from

the partially crystalline 1995 magma and possibly from

other parts of the magma storage region.

Magma volumes and storage architecture

The distinctive chemical signatures of groundmass glasses

from individual Ruapehu eruptions suggest that each

eruption tapped a slightly different magma. We know from

field investigations and measurements of the eruptive

deposits that the eruptive volumes were very small

(between 1945 and 1996). The total volume of magma

erupted between 1945 and 1996 is approximately

3.6 9 107 m3. This total magma volume estimate is dom-

inated by the 1995–1996 eruptions and is very small in

comparison with a single, moderately sized andesitic

eruption (e.g., Bezymianny; Belousov et al. 2002, Colima;

Saucedo et al. 2010).

Based on the saturation pressure calculated from the

H2O and CO2 content of phenocryst-hosted melt inclu-

sions, we determined a magma storage depth of *2–9 km.

This compares well to magnetotelluric (MT) data (Ingham

et al. 2009) and seismic tomography (Rowlands et al. 2005)

from Ruapehu. Ingham et al. (2009) observed a diffuse and

weak low-resistivity anomaly that extends to *6 km,

which they interpreted to be a dyke system. From *6 to

more than 10 km and slightly east of the cone, a more

intense low-resistivity anomaly (melt-bearing zone) was

identified using both 2-D and 3-D inversions. The seismic

tomography data also show a low velocity zone to the east

of the cone from *3 to *9 km depth, although these data

have been interpreted as a combination of crustal down-

warping and the presence of thick volcaniclastic sediments

(Rowlands et al. 2005). Our data are consistent with a

magma storage region (possibly in the form of discrete, yet

closely spaced sills and dykes) down to 9 or 10 km

(Fig. 11). We suggest that prior to eruption, high-angle (c.

80) dykes (calculated from geophysical data) transport

magma to the active vent beneath Crater Lake. These dykes

probably pass through and interact with partial melt zones

that may take the form of small sills (crystal mush zones)

(Fig. 11). The sill-like nature of these bodies enables dif-

ferent eruptions to interact with mushes that were chemi-

cally and physically isolated from one another, as evinced

by the melt inclusion data discussed above.

The small volumes of these magma bodies beneath

Ruapehu are unlikely to be readily imaged by geophysical

techniques such as MT or seismic tomography. In fact,

Ingham et al. (2009) use their MT data to suggest that it is

unlikely that large volume magma bodies are able to

accumulate in the shallow crust beneath Ruapehu.

Although physically isolated at shallow depths, the

various Ruapehu magmas are likely to be genetically

linked at depth, and furthermore, it is possible that a tem-

poral trend can be drawn based on our groundmass glass

and major element chemistry data, whereby eruptions have

become more mafic with time (since 1945). However,

given that there are a number of eruptions for which we do

not have samples or analyses (including 1975), we can only

speculate that a temporal-chemical trend exists.

Conclusions

Historical eruptions at Ruapehu (1945, 1969, 1971, 1977,

1995, and 1996) are characterised by very small volume

magmas, each with a unique chemical composition and

history. Volatile contents of melt inclusions and crystal-

melt barometry have constrained the depth at which these

magmas originated to be *2 to *9 km, which corre-

sponds well to geophysical data. These small volume melts

probably resided as distinct and closely spaced sills or

dykes from 2 to 9 km. Before an eruption, magma was

injected into the sill/dyke system leading to common

magma-mush and magma–magma interaction. Most mag-

mas interacted with crystal mush zones (at\*3 km depth)

formed from antecedent magmas during ascent and even-

tually eruption. Due to their small volumes, Ruapehu

magmas since 1945 show sensitivity to interaction between

magmas and crystal mush zones that would be difficult to

determine in larger magmatic systems. Therefore, data

from Ruapehu offer a unique insight into the small-scale

interactions that magmas experience on their ascent to

eruption.

We have shown that the chemical composition of phe-

nocryst-hosted melt inclusions is often distinct from the

groundmass glass. This implies that a significant proportion

of the crystals are antecrysts; in some cases, antecrysts

have incorporated rare melt inclusions from the new melt.

Contrib Mineral Petrol

123

In that respect, it is clear that the interpretation of mag-

matic processes at depth can only be achieved from cog-

nate, rather than exotic, melt inclusions.

Ruapehu magmas are low in H2O but CO2-rich com-

pared to intermediate magmas from subduction settings

elsewhere. This relatively low concentration of volatiles

and the very small volumes of magma combine to account

for the low explosivity and short duration of most eruptions

at Ruapehu. While the largest eruptive episode (in

1995–1996) produced plumes to 20 km, the volatile con-

tent is similar to the smallest episode analysed (in 1971).

Therefore, the controls on the size of eruptions at Ruapehu

are not determined by the volatile content of magma alone.

Acknowledgments This work was funded by the New Zealand

Ministry of Science and Innovation (MSI) Geological Hazards Pro-

gramme (GHZ) in the form of a PhD studentship to GK at the Uni-

versity of Bristol. Holly Goddard and Neville Orr are thanked for their

assistance with sample preparation. We gratefully acknowledge sup-

port from NERC for access to the SIMS facility, Edinburgh, where

Cees-Jan de Hoog provided expert guidance and patience. Stuart

Kearns and Ben Buse are thanked for their support with EPMA and

SEM analyses. JB is supported by ERC Advanced Grant ‘‘CRIT-

MAG’’ and a Royal Society Wolfson Research Merit Award. KC

acknowledges funding from the AXA Research Fund. Two anony-

mous reviewers are thanked for constructive and helpful reviews that

helped us clarify and significantly improve the manuscript.

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